Salty Matters

The Blog is written by me, John Warren. Once every three or four weeks or so I will post an article or two on an evaporite topic that has piqued my interest. On the Saltwork Publications webpage (under "the Works") there is a growing library of pdfs and epubs based on these blogs. These articles on the website have much higher resolution extractable graphics in than in the blog. There is also a link to this set of pdfs and epubs on the home page (www.saltworkconsultants.com).

Lapis Lazuli: A metamorphosed evaporite

John Warren - Friday, November 13, 2015

Introduction 

Precious stones and [1]gems are rare by definition; hence need exceptional geologic conditions to give rise to gem-quality materials. A nexus across most natural gem-forming environments is the requirement for hydrous typically saline to hypersaline solutions, apt to precipitate euhedral crystals in a void or a pressure shadow, from fluids that contain elevated and unusual levels of particular constituents, including chromophores; hence pegmatites, volcanics and meta-evaporites are commonplace hosts for natural gemstones. Fluids promoting the growth of gem-quality crystals typically include the availability of uncommon major constituents, along with the presence of adequate chromophores [2] , as well limited concentrations of undesirable elements. Another need is open fluid space in an environment conducive to growing crystals of sufficient size and transparency. This general statement of requirements to form a precious stone also encapsulates why some gems have meta-evaporitic associations.

We know that depositional units of evaporite salts typically disappear or transform into other mineral phases by the early greenschist phase (Warren, 2016). As this happens the dissolution/transformation releases a pulse of hot basinal chloride waters that can carry gold and base metals (topics for another blog) as well as leaching and carrying elements such as beryllium, chromium and vanadium (chromophores) from adjacent organic-rich shales. Trace elements also tend to be enriched in the more evolved depositional brines that precipitate in later minerals in a primary evaporite precipitative series. Later, at the same time as halite dissolves or transforms on entry into the metamorphic realm, anhydrite layers and masses typically remain into the more evolved portions of the metamorphic realm (amphibolite-granulite facies). The volume loss associated with the dissolution/transformation of meta-evaporites facilitates the formation of open fluid space (sometimes pressurized) in veins and fractures so favoring sites that then allow the free growth of precious stones and euhedral ruby, tourmaline and emerald gemstones.

I am not saying all precious stones and gems are associated with evaporites, many natural gemstone settings are not, but the lapis lazuli of Afghanistan, the melon-sized perfect rubies of Myanmar, the prolific emerald fields in Columbia, and the tsavorite deposits of east Africa likely are (Garnier et al., 2008; Giuliani et al 2005; Feneyrol et al., 2013). In this article I will focus on lapis lazuli, for a discussion of other semi-precious stones and gems that are meta-evaporites see Chapter 14 in Warren (2016).


 

Lapis lazuli

Lapis Lazuli is the metamorphic remnant of a sodic-rich, quartz-absent, evaporite mineral assemblage. It is composed of an accumulation of minerals not a single mineral (unlike rubies and emeralds); it is mostly lazurite (Na, Ca)8(AlSiO4)6(S, SO4,Cl)1-2), typically at levels of 30-40%. Lapis gemstone also contains calcite (white veins), sodalite (blue), and pyrite (gold flecks of color). Dependent on metamorphic history and protolith chemistry, other common minerals in lapis include; augite, diopside, enstatite, mica, haüyanite, hornblende, and nosean. Some specimens also contain trace amounts of the sulphur-rich mineral lollingite (var. geyerite). Lazurite is a member of the sodalite group of feldspathoid minerals (Table 1). Feldspathoids have chemistries that are close to those of the alkali feldspars, but are poor in silica. If free quartz were present at the time of formation it would have reacted with any feldspathoid precursor to form feldspar not lazurite. Natural lazurite contains both sulphide and sulphate sulphur, in addition to calcium and sodium, and so is sometimes classified as a sulphide-bearing haüyne (Figure 1). Sulphur gives lazurite its characteristically intense blue color, which comes from three polysulphide units made up of three sulphur atoms having a single negative charge. The S3- ion in the sulphur has a total of 19 electrons in molecular orbitals and a transition among these orbitals produces a strong absorption band at 600 nm, giving a strong blue color, with yellow overtones. The intensity of the gem’s blue is increased with increasing sulphur and calcium content, while a green color is the result of insufficient sulphur (O’Donoghue, 2006, p. 329).

 

Other members of the sodalite group include sodalite and nosean (Table 1). Sodalite is the most sodium-rich member of the sodalite group and differs from the other minerals of the group in that its lattice retains chlorine. Interestingly, sodalite can be created in the laboratory by heating muscovite or kaolinite in the presence of NaCl at temperatures of 500°C or more. In the literature, the commonly accepted origin of lazurite is through contact metamorphism and metasomatism of dolomitic limestone. Such a metasedimentary system also requires a source of sodium, chlorine and sulphur; the obvious source is interbedded evaporites in the protolith, as is seen in plots of its molecular constituents (Aleksandrov and Senin, 2006).


Lapis lazuli from the Precambrian of Baffin Island, Canada (Figure 1), and from Edwards, New York, are meta-evaporites with evaporite remnants (anhydrites) remaining in the same series, as are the lapis lazuli deposits at Sar-e-Sang in the Kokcha valley, Afghanistan and the lapis deposits in Liadjuar-Dara region (“River of Lazurite”) at an altitude of 5000 m in the Pamir Mountains, Tajikistan (Webster, 1975). Throughout history its bright blue color has made lapis, mostly from Sar-e-Sang, a valued gem commodity. First mined 6000 years ago, the Sar-e-Sang lapis was transported to Egypt and present day Iraq and later to Europe where it was used in jewelry and for ornamental stone[3]. Europeans even ground down the rock into an expensive powdered pigment for paints called “ultramarine”.

Lapis deposits in Lake Harbor on Baffin Island and in the Edwards Mine, New York, were produced by high-grade metamorphism of a sulphate-halite-marble protolith (Hogarth and Griffin, 1978). The anhydrites preserved near Balmat are remnants of this sequence. On Baffin Island the two main lapis lazuli lenses, some 1.6 km apart, lie at the structural top of two sequences of dolomitic marble, the thicker lens being approximately 150 m across (Figure 1b). The elongation of both lenses parallels the local layering and foliation and shows a well-developed layering parallel to the regional foliation, giving additional evidence of its sedimentary protolith to the deposits. The Main and Northern bodies constitute diopside–lazurite rocks of variable gem quality and are localized in marbles among biotite gneisses. The Main (Southern) occurrence is as long as 170 m and 6 m thick. In these deposits, sheets of high-quality lazurite (up to 1 m thick) contain variable amounts of relict diopside and plagioclase, as well as newly formed haüyne, nepheline, or phlogopite. The quantitative proportions of these minerals define the color of the rock, which changes to a more intense blue with heating. The Northern occurrence (25×36 m in size) is less rich than the Main occurrence and consists of small (no more than 1 m) lenses showing disseminated lazurite, which imparts a bluish green color to the polished surface of the rock. Chlorine and sulphur in the various lazurites, accessory pyrite, and pyrrhotite were derived from metamorphosed gypsum-, anhydrite-, and evaporitic-carbonate protoliths (Hogarth and Griffin, 1978).


In the Lake Baikal lazurite occurrences, there is once again a strong association between marble of the Perval’na Group and lazurite occurrence (Figure 2a). For example, the Slyudyanka deposit is hosted in diopside skarns and spinel–forsterite calciphyres, developed from metamorphically-evolved evaporitic dolomites (Aleksandrov and Senin, 2006). The Slyudyanka deposit shows clearly pronounced metasomatic zoning, which was associated with the prograde magnesian skarn stage and was overprinted by retrograde postmagmatic assemblages, that formed together with lazurite-bearing rocks under the influence of saline alkaline S–Cl-bearing hydrothermal solutions. These solutions also caused microclinization of blocks of leucocratic granite with the formation of lazurite in the some of the inner skarn zones. Potassium solutions caused phlogopitization of the host rocks.

Likewise, scapolite and magnesian whiteschists are typically saline mineral phases in the classic deposits of the Sar-e-Sang District (Figure 2b; Faryad, 2002). There, the lapis is composed of a combination of lazurite, diopside, calcite and pyrite and occurs in beds and lenses up to 4 meters thick within a scapolitic magnesian-marble skarn near the center of the Hindu Kush granitic massif. It is typically interlayered with, or forms veins and lenses within a gneissic and pegmatitic host. Lens-shaped lodes are typically hosted in orthoclase–microcline–perthite hornfels containing albite and quartz (Figure 2b). Lazurite bodies at the Sar-e-Sang deposit are associated with diopside metasomatites bearing nepheline, pale blue haüyne, and blue lazurite, and some lazurite-rich zones can contain up to 40-90 vol% lazurite. The rocks also contain diopside, haüyne, afghanite, and nepheline, as well as disseminated pyrite replaced by pyrrhotite. Pockets of near pure lapis lazuli can be up to 40m across and occasionally up to a meter.

Lapis lazuli in the North Italian Mountains of Colorado occurs in impure marbles in a meta-evaporitic skarn near the contact with the Eocene-age quartz monzonite and quartz diorites of the Italian Mountain stocks (Hogarth and Griffin, 1980; Mauger, 2007). There, near vertical Pennsylvanian black shales and carbonates along the west margin of the intrusive have been converted to phlogopite-diopside-andalusite hornfels and scapolite-diopside skarns with minor analcime. Compared to Sar-e-Sang, lapis in this skarn deposit is of inferior quality. It forms as deep blue lazurite granules in fine-grained forsterite-Ti phlogopite-calcite skarns and calcite marbles with diopside, Ti-phlogopite and pyrite. The hosting sediments (Mississippian limestones and Devonian sandstones) define along the NE margin of the pluton, while the NaCl came from dissolution of once nearby halite or dissolution-derived saline surface waters and shallow groundwaters moving south from the Eagle Basin.

High quality lapis is also mined from a limestone-granite skarn contact in the Chilean Andes (3500 m elevation) in the headwaters of the Cazadero and Vias River, Ovalle, Coquimbo, Chile. The lapis there is good quality, although somewhat paler than Sar-e-Sang and, like the Baikal lapis deposits of Russia, is associated with wollastonite not diopside, making it a less attractive gem. The Chilean lapis occurs in an association of phlogopite, sodalite, calcite and pyrite (Coenraads and Canut de Bon, 2000).

Meta-evaporites in the Sar-e-Sang region of Afghanistan exhibit mosaic equilibria across small volumes (in the cm3 range) within a talc-kyanite schist (whiteschist) host. The microscale mineral variations are characterized by variations in mineral assemblages conventionally attributed to vastly different pressure/temperature conditions during regional metamorphism.

On the basis of petrographic and microprobe studies, these assemblages are attributed to three consecutive stages of metamorphism of a chemically exceptional rock with a composition that falls largely into the model system MgO-Al2O3-SiO2-H2O (Figure 3; Schreyer and Abraham, 1976). Stage 1, typified by Mg chlorite-quartz -talc and some paragonite, was followed during stage 2 by talc-kyanite, Mg [4]gedrite-quartz and the growth of large dravites (magnesian tourmalines). Microprobe analyses of the phases, gedrite and talc, indicate variable degrees of sodium incorporation into these phases according to the substitution NaAl—>Si. In stage 3, pure Mg cordierite formed with or without corundum and/or talc, and the kyanite was partly converted into sillimanite. Pressure and temperature during this final stage of metamorphism was near 5-6 kb and 640°C.


Schreyer and Abraham (1976) concluded that chemical variations in the metamorphic fluids were generated by progressive metamorphism and mobilization of an evaporite deposit. Relict anhydrite and gypsum(rehydrated anhydrite) still occur in the Sar-e-Sang area. Whiteschists and the associated lapis lazuli deposits of the region are part of a highly metamorphosed evaporitic succession. Salts have largely vanished due to ongoing melting and volatilizations. The preservation of the three stage succession of mineral assemblages, across such small scales and yet related to each other through isochemical reactions, means that the main factors governing the metamorphic history of this whiteschist were compositional changes of the coexisting fluids with time. Under this scenario any pressure-temperature variations were subordinate and the chemistry of the fluids evolved as the evaporites underwent metasomatic alteration.

The sedimentary pelitic layers of this precursor evaporitic sequence first underwent a period of metamorphism in which fluid pressures approached lithostatic (stage 1). Subsequently at higher metamorphic grades, with the beginning of mobilization of the salts, the metamorphic fluids became increasingly enriched in ions such as Na+, Mg2+, Cl-, SO42-, BO33-, etc., so that water fugacity dropped considerably. This period is represented by stage 2 of the whiteschist metamorphism and was characterized by strong metasomatism that led, for example, to the growth of dravite and the amphibolite, gedrite. The physical and chemical character of stage 3 is less clearly defined. Kyanite/sillimanite inversion requires an increase in temperature or a decrease in pressure, or both; but changes in the composition of a coexisting gas phase may have played an additional role in the formation of cordierite.

Unlike classic metamorphic associations, the meta-evaporite-derived assemblage in Afghanistan may in a single thin section entrain mineral assemblages that conventionally would be assigned to the greenschist facies, the hornfels facies, and to a high pressure (amphibolite) regime. The assemblages are in effect mosaic equilibria that reflect changes in fluid composition generated from a metamorphosing evaporite pile over time and only to a lesser degree by regional evolution of total temperature and pressure. Once again, evaporites generate unusual responses compared to the general responses of metasediments.

In a refinement paper discussing the likely relationships between evaporites and whiteschists, Franz et al., 2013 note that whiteschist mineral assemblages are stable up to pressures of more than 4 GPa, but may already form at pressures of 0.5 GPa. Their formation largely depends on the composition of the protolith and requires elevated contents of Al and Mg as well as low Fe, Ca, and Na contents, as otherwise chloritoid, amphibole, feldspar, or omphacite are formed instead of kyanite or talc. They go on to note that the stability field of a whiteschist mineral assemblage strongly depends on XCO2 and fO2: at low values of XCO2, CO2 binds Mg to carbonates strongly reducing the whiteschist stability field, whereas high fO2 enlarges the stability field and stabilizes yoderite [Mg(Al,Fe3+)3(SiO4)2O(OH)].

The scarcity of whiteschist is not necessarily due to unusual P–T conditions, but to the restricted range of suitable protolith compositions and the spatial distribution of these protoliths: (1) continental sedimentary rocks and (2) hydrothermally and metasomatically altered felsic to mafic rocks. They argue continental sedimentary rocks that may produce whiteschist mineral assemblages typically have been deposited under arid climatic conditions in closed evaporite basins and may be restricted to relatively low latitudes. These rocks typically contain large amounts of palygorskite and sepiolite. Franz et al., (2013) conclude whiteschist assemblages typically are only found in settings of continental collision or where continental lacustrine fragments were involved in subduction.

In my opinion, the mosaic signature of the precursor mineral phases in the typical Sar-e-San lapis lazuli is a metamorphically-evolved response to the combination of precursor permeability and stability contrasts typical of variably-cemented halite mosaic sediments in what were likely haloturbated and variably cemented saline continental lacustrine precursors.

References

Aleksandrov, S., and V. Senin, 2006, Genesis and composition of lazurite in magnesian skarns: Geochemistry International, v. 44, p. 976-988.

Coenraads, R., and C. C. de Bon, 2000, Lapis Lazuli from the Coquimbo Region, Chile: Gems & Gemology, v. 36, p. 28-41.

Faryad, S. W., 2002, Metamorphic Conditions and Fluid Compositions of Scapolite-Bearing Rocks from the Lapis Lazuli Deposit at Sare Sang, Afghanistan: Journal of Petrology, v. 43, p. 725-747.

Feneyrol, J., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. E. Martelat, P. Monié, J. Dubessy, C. Rollion-Bard, E. Le Goff, E. Malisa, A. F. M. Rakotondrazafy, V. Pardieu, T. Kahn, D. Ichang'i, E. Venance, N. R. Voarintsoa, M. M. Ranatsenho, C. Simonet, E. Omito, C. Nyamai, and M. Saul, 2013, New aspects and perspectives on tsavorite deposits: Ore Geology Reviews, v. 53, p. 1-25.

Franz, L., R. L. Romer, and C. Capitani, 2013, Protoliths and phase petrology of whiteschists: Contributions to Mineralogy and Petrology, v. 166, p. 255-274.

Garnier, V., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. Dubessy, D. Banks, H. Q. Vinh, T. Lhomme, H. Maluski, A. Pecher, K. A. Bakhsh, P. Van Long, P. T. Trinh, and D. Schwarz, 2008, Marble-hosted ruby deposits from Central and Southeast Asia: Towards a new genetic model: Ore Geology Reviews, v. 34, p. 169-191.

Giuliani, G., A. E. Fallick, V. Garnier, C. France-Lanord, D. Ohnenstetter, and D. Schwarz, 2005, Oxygen isotope composition as a tracer for the origins of rubies and sapphires: Geology, v. 33, p. 249-252.

Hogarth, D. D., and W. L. Griffin, 1978, Lapis lazuli from Baffin Island; a Precambrian meta-evaporite: Lithos, v. 11, p. 37-60.

Mauger, R. L., 2007, Contact metamorphism-metasomatism associated with the latest Eocene northern Italian Mountain granite intrusion, Gunnison County, Colorado: Abstracts with Programs - Geological Society of America, v. 39, p. 394.

O'Donoghue, M., 2006, Gems; Their Sources, Descriptions and Identification (6th Edition): Amsterdam, Elsevier, 873 p.

Schreyer, W., and K. Abraham, 1976, Three-stage metamorphic history of a whiteschist from Sar e Sang, Afghanistan, as part of a former evaporite deposit: Contributions to Mineralogy & Petrology, v. 59, p. 111-130.

Von Rosen, L., 1990, Lapis lazuli in archaelogical contexts, in P. Aströms, ed., Studies in Mediterranean Archaeology and Literature, v. 93, Partille, Sweden.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Dec. 2015: Berlin, Springer, 1600 p.

Webster, R., 1975, Gems, their sources, descriptions and identification: London, Newnes, Butterworths.

Wood, J., 1841, John Wood. A Personal Narrative of a Journey to the Source of the River Oxus by the Route of the Indus, Kabul, and Badakhshan, Performed under the Sanction of the Supreme Government of India, in the Years 1836, 1837, and 1838. Avaialble as Elibron Classics, 2001, 458 pages. Replica of 1841 edition by John Murray, London.



[1] The Ancient Greeks, distinguished between precious and semi-precious stones; similar distinctions were made in other ancient cultures. In modern usage, the precious stones are diamond, ruby, sapphire and emerald, with all other gemstones, including lapis lazuli, being semi-precious.

[2] Chromophore; is the part of a gem lattice responsible for its color. A gem’s color arises when a molecule absorbs certain wavelengths of visible light and transmits or reflects others. It is a structural feature in the lattice indicative of the presence of a gem-specific electron configuration of the ions in its crystal lattice; such as transition metal ions (Cr, V or Fe) occupying several different coordination sites. For example, ferrous iron (Fe2+) or ferric iron (Fe3+), the ferrous ion in peridot causes the green color and ferric ion causes the yellow color in chrysoberyl. This color effect has important uses in heat-treatment gemstones such as blue color in heat-treated sapphire.

[3] Lapis is the Latin word for “stone” and lazuli is the genitive form of the Medieval Latin lazulum meaning blue, which was taken originally from the Persian lāžaward, the name of a place where lapis lazuli was mined. Taken as a whole, lapis lazuli originally meant “stone of Lāzhward.” With time, the name of the place came to be associated with the stone mined there and, eventually, with its bright blue color.

Lapis lazuli’s use as jewelery can be traced back to the 5th millennium B.C.E. with the discovery of beads at a cemetery outside the temple walls of Eridu (Sumer) in southern Babylonia (Von Rosen, 1990) and was used as glyptic from then until now in the manufacture of jewels, amulets, seals and inlays. To the ancient Egyptians, it was considered a gem representing the skies or heaven, thus was thought to denote light, truth and wisdom. It was thus often shaped into eye-shaped gems and was worn by judges in ancient Egypt. A lapis amulet graced the brow of Ra. Lapis is noted in Revelations in Christian mythology as a stone in the Breastplate of Aaron. In China, lapis was worn during the Manchu dynasty for services in the Temple of Heaven. The Romans and Greeks used it as a cure for fever and melancholy. It also glazed the bricks that formed the spectacularly blue “Gate of Kings” or Ishtar entryway to the ancient city of Babylon (≈1800 BC).

The Sar-e-Sang region has supplied much of the gem quality lapis to the world. One of the first European explorers to the region (Wood, 1841) described mining methods in use at that time. Camel-thorn and tamarisk twigs were collected from the valley below and carried up the steep path to the mine. When sufficient fuel had been collected, it was piled against the rock face and a fire was lit. When the rock was hot, cold water, which also had to be carried up the steep 350 m ascent from the valley floor, was thrown onto it. The rock cracked and split, enabling further work to be done with the primitive tools available (pick, hammer and chisel) in order to extract the lapis lazuli from its marble host rock.

[4] Gedrite is a silicate mineral of the amphibole group with formula: (Mg;Fe2+)2[(Mg;Fe2+)3Al2](Si6Al2)O22(OH)2

Saline Clays

John Warren - Thursday, July 23, 2015

When discussing evaporites we typically focus on the formation and alteration of the various evaporite salts and their diagenetic evolution, but the same evolving saline hydrologies can also drive the formation and alteration of clays (Table 1). Many authigenic clay minerals formed in hypersaline settings are enriched in magnesium (Fisher, 1988), but authigenic clays do not make up the greater volumes of clay in modern or ancient salt lakes. Most of the clays in salt lakes and playas are detrital and reflect compositions of older argillaceous formations in the palaeodrainage areas. Illite, kaolinite, chlorite, dioctahedral smectite and a number of mixed-layers clays are commonplace detrital clay minerals in saline formations (Figure 1; Calvo et al., 1999). Widespread flocculation of clays is an effective sedimenter of suspended clay wherever freshwater runoff and streams flood an area of standing saline water. Thus the composition of initial clay sediments in a playa largely reflects that of the minerals carried as suspended load into the lacustrine depression.


The magnitude of detrital clastic input is thought to be a significant factor in the relative volume of authigenic clay. Regions with rapid deposition of clays, tied to high detrital inputs, tend to be areas where the authigenic clay component is swamped by the high detrital input. Clay authigenesis in evaporitic basins is favoured in marginal playa areas where rates of detrital clay input are low (Figure 1). This encompasses interdunal depressions, peripheral sandflats and muddy carbonate flats. In these low sedimentation areas the transformation of precursor clays is more effective, driven by episode surface inflow and groundwater discharge (Calvo et al., 1999). Highly reactive nearsurface and surface conditions are favoured by inherently large variations in pore water salinity, pH and pCO2 levels.

 

Clay authigenesis in many saline depressions is driven by pedogenesis, especially in the marginal areas where sedimentation rates are low and subaerial exposure dominates at the sedimentation surface. Below the surface episodic wet-dry cycles means neoformed clays are the byproduct of complex reactions between Na and Mg-rich interstitial brines and detrital silicates. Pedogenic processes account for the formation of widespread lake margin palygorskite and sepiolite, typically in association with the creation of calcretes, dolocretes and silcretes. In cases where palygorskite dominates the soil profile, they are sometimes described as palycretes. Zeolites can also form from saline groundwaters in saline lake-margin pedogenic settings (Figure 2). Artesian and phreatic groundwater discharge through springs into the lake margin areas also plays a significant role in the formation of other authigenic clays, as in saline lakes at the foot of Mt Kilimanjaro, in Tanzania and Kenya (Hay et al., 1995).

 

Hypersaline brines in modern, marine-edge evaporite basins can also enhance clay authigenesis even in settings where thermal and saline stresses keep both organic and inorganic carbon concentrations in the sediments unusually low relative to coastal marine environments with lower salinities (Martini et al., 2002). This is the case in Salina Ometepec where sediment pore waters exhibit little microbial sulphate reduction, and dissolved inorganic C contents are also very low. Instead of carbonate alteration (dolomitisation) in the Mg brine, authigenic K-rich Mg-smectite (saponite) formation is occurring, driven by the concurrent processes of brine concentration, selective dissolution of K- and Mg-bearing salts, and dissolution of detrital aluminosilicates. Salina Ometepec pore waters at a depth of 1 m have 87Sr/86Sr ratios that require input of Sr that is less radiogenic than that of Gulf of California seawater. This Sr is likely derived from weathering and leaching of detrital aluminosilicates from nearby volcaniclastic sources. Although rare in Holocene successions, similar Mg-rich authigenic clay assemblages are well documented in Palaeozoic evaporite basins (Bodine, 1983; Janks et al., 1992; Andreason,1992).

Once precipitated in an evaporite basin, authigenic clays can be retransported further out into the saline depression and in more humid climatic stages may even end up on the floor of freshwater lakes (Figure 1). This situation is seen in lacustrine sequences from the Miocene formations of the Madrid Basin (Bellanca et al., 1992) where significant amounts of palygorskite and sepiolite occur as either mud chips or clay aggregates in the basal part of a fresher water lacustrine unit. Eolian transport of saltating clay pellets or dust suspensions may also contribute to the transport of authigenic clays from marginal to more central areas. This sometimes leads to problems of interpretation of detrital versus authigenic in ancient lacustrine successions subject to oscillations in climate, especially when detrital clays are partially or fully inherited from arid soils.

Sepiolite, interstratified Mg-Smectite and palygorskite form authigenic phases in the Quaternary sediments of the Double Lakes Formation, Texas (Webster and Jones, 1994). The dominance of each of these minerals in separate horizons represents evaporative shifts in salinity at the time they precipitated. Sepiolite is thought to indicate a brackish lake, while Mg-smectite indicates more saline conditions. Palygorskite is interpreted as a saline pore water precipitate in the arid soils of the playa stage. Likewise Jones (1986) interpreted authigenic Mg-smectites (e.g. stevensite) as requiring higher salinity than sepiolite. Mg-silicates also define saline lake clays in Great Salt Lake (Spencer, 1983) and some Bolivian salars (Badaut and Risacher, 1983). In Bolivia, the authigenic Mg-smectite replaces the biogenic silica in diatom frustules and requires a pH in excess of 8.2. Authigenic stevensite occurs in unconsolidated muds underlying saline crusts in the interdunal depressions of northern Lake Chad and as small aragonite-associated oolites on the lake floor (Gac 1980, Darragi and Tardy, 1987). Similar stevensite oolites have been found in the Eocene Green River lacustrine basin. Stevensite is also an early authigenic phase in the modern carbonate thrombolites in the hyposaline Lake Clifton, Australia (Burne et al., 2014). Authigenic sepiolite associated with calcite, gypsum and dolomite occurs about the margin of Saline Valley Playa, California and the edges of saline pans in the Kalahari of southern Africa (Hardie, 1968; Kautz and Porada, 1976). Palygorskite, sepiolite and authigenic smectite are commonplace precipitates in calcretes of groundwater discharge playas in inland Australia (Arakel et al., 1990).

Clearly, palygorskite and sepiolite (both two-chain structure fibrous clays) occur worldwide as authigenic phases in the soils and palaeosols of arid and semi-arid regions, but the mode of precipitation is still not well understood (Singer, 1979). Both minerals are common in environments with elevated levels of magnesium and silica. Hence they form in alkaline lakes and caliche, as well as in deep sea sediments and Hydrothermal alteration products; Folk and Rasbury (2007) argue there may also be a microbial association to their formation, at least in some Texan caliches. Jones (1986) concluded sepiolite in the calcic soils of southwest Nevada required percolation of high salinity groundwaters. Magnesium and silica solutes were supplied by the weathering of nearby pyroclastics and carbonates. Sepiolite has replaced magnesite pebbles, from the edges in, during freshened highstand intervals in Miocene Lake Eskisehir in Turkey (Ece, 1998). Palygorskite in calcic soils is thought to be the result of incongruent dissolution of pre-existing clays (Jones and Galán, 1988). Fibrous clays degrade when climate becomes more humid and alter to smectite. Paquet and Millot (1972) conclude that the transformation takes place when mean rainfall exceeds 300 mm and Calvo et al. (1999) suggested the transformation can be used as a palaeoclimatic indicator.

Alunite (KAl3(SO4)2(OH)6) is a common clay product in acid saline lacustrine settings, but can also form diagenetically in regions where sulphate reduction is occurring. It is thought to be derived by the reaction of clay minerals with sulphuric acid created by oxidation of sulphides or H2S at a redox boundary. It is a common product where clays are present in zones of sulphate reduction and examples have been documented in the Middle Miocene gypsums of the Gulf of Suez (Rouchy et al., 1995) and the Upper Miocene gypsums of the Lorca Basin in Spain (Rouchy et al., 1998).

Even the smectite to illite transformation, which is used as an indicator of diagenetic intensity and clay transformations occurring at higher temperatures may be influenced by salinity. This makes illite crystallinity a less reliable indicator of diagenetic stage in environments with saline pore fluids (Honty et al., 2004). Turner and Fishman (1991) found illite-smectite mixed layer clays having a range of expandabilities in altered tuff beds in a Jurassic lake in the Morrison Formation (Eastern Colorado Plateau, USA). The observed clays did not experience deep burial, and did not undergo hydrothermal alteration. The illite content generally increases from the lake margin (100–70% smectite) to the lake centre (30–0% smectite) and follows a lateral hydrogeochemical gradient, which was characterized by increasing salinity and alkalinity (Figure 3). It seems that in a saline depositional setting, solution chemistry is a principal factor controlling the smectite to illite proportion. Illite-smectite can form from smectite at low temperatures in several ways (see Honty et al., 2004), but forms best in saline environments subject to wetting and drying cycles, which is a hydrology exemplified in salt lakes and playas. In the presence of K+ ions, alternating wetting and drying leads to irreversible fixation of K and the formation of illite layers. Illite-smectite clays forming at elevated pH may not even require wetting and drying cycles.


References

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